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    (a) Climatological-mean winds at 850 hPa and (b) variance of 20–80-day-filtered OLR anomalies during June–September for the period 1979–98. The dashed box in (a) and (b) represents the major monsoon trough region (12°–22°N, 123°–133°E) for an intensity index, and areas more than 400 (W m−2)2 in (b) are shaded. The composite wind at 850 hPa during (c) the active phase and (d) the inactive phase. The dashed box in (c) and (d) represents the area (10°–30°N, 112°–132°E) in Fig. 2.

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    The vertical profiles of area-averaged 20° × 20° (a) tangential wind, (b) temperature, (c) divergence, and (d) specific humidity during the active (solid) and inactive (dashed) phases of monsoon trough ISO centered in the circulation anomaly. The vertical dashed line indicates the zero value of each parameter.

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    Wind fields (vectors) and total wind speed (shaded) at 850 hPa in (a)–(c) CTL, (d)–(f) ACV relative to AC, and (g)–(i) IACV relative to IAC at t = (left) 0, (center) 24, and (right) 48 h from 27-km simulation data.

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    Time evolution of (a) the MSLP and (b) the MAMW speed at 10 m for CTL (black solid), ACV (red dashed), and IACV (blue dotted). The abscissa represents time (h) and the ordinate corresponds to the intensity.

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    The radial distribution of the (a)–(d) azimuthal-mean perturbation radial wind, (e)–(h) tangential wind, and (i)–(l) geopotential at 1000 hPa from t = 1 to 7 h with an interval of 2 h for ACV relative to AC (black) and IACV relative to IAC (red).

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    The vertical–radial cross section of azimuthal-mean perturbation radial wind (m s−1) from t = 1 to 7 h with an interval of 2 h for (a)–(d) ACV relative to AC and for (e)–(h) IACV relative to IAC and (i)–(l) CTL.

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    The vertical–radial cross section of azimuthal-mean (m s−2) at t = 1 h between the (a) active phase and (b) inactive phase.

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    The vertical–radial cross section of (a)–(c) azimuthal-mean perturbation radial wind (m s−1), (d)–(f) divergence (10−5 s−1), (g)–(i) vertical velocity (10−3 m s−1), and (j)–(l) diabatic heating (10−4 K s−1) at (left) t = 5.5, (center) 6, and (right) 6.5 h for ACV relative to AC.

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    As in Fig. 8, but for IACV relative to IAC.

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    Vertical–time cross section of (a),(b) perturbation diabatic heating (10−4 K s−1) and (c),(d) perturbation vertical velocity (10−2 m s−1) averaged over 420 km × 420 km centered on the TC center in the first 10 h for (left) ACV relative to AC and (right) IACV relative to IAC.

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    Time (24-h running mean) evolutions of radial-mean (0–180 km) (a) radial wind at 10 m, (b) tangential wind at 10 m, (c) vertically integrated (1000–200 hPa) diabatic heating, and (d) vertically integrated (1000–200 hPa) vertical velocity for ACV relative to AC (red) and IACV relative to IAC (blue) cases.

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    The vertical–radial cross section of azimuthal-mean (a)–(f) vertical velocity (10−2 m s−1) and (g)–(l) diabatic heating (10−4 K s−1) fields for (a)–(c),(g)–(i) ACV relative to AC and (d)–(f),(j)–(l) IACV relative to IAC cases, averaged at 0–24, 24–48, and 48–72 h.

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    As in Fig. 4, but for CTL (black solid), ACV (red dashed), ACV_SH (blue dotted), and ACV_NOSH (green dash–dotted).

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    The vertical–radial cross section of azimuthal-mean (a)–(c) perturbation radial wind (m s−1), (d)–(f) divergence (10−5 s−1), (g)–(i) vertical velocity (10−3 m s−1), and (j)–(l) diabatic heating (10−4 K s−1) at (left) t = 5, (center) 7, and (right) 9 h for CTL run.

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    As in Fig. 14, but for ACV_NOSH relative to AC_NOSH.

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    As in Fig. 14, but for ACV_SH relative to AC_SH.

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    The vertical–radial cross section of (a),(c),(e) perturbation vertical velocity and (b),(d),(f) perturbation diabatic heating averaged over 420 km × 420 km centered on the TC center in the first 10 h for (a),(b) CTL, (c),(d) ACV_NOSH relative to AC_NOSH, and (e),(f) ACV_SH relative to AC_SH.

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    As in Fig. 4, but for CTL (black solid), IACV (red dashed), IACV_SH (blue dotted), and IACV_NOSH (green dash–dotted).

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    As in Fig. 4, but for CTL (black solid), ACV (red dashed), IACV (blue dotted), ACV_SH + IACV_NOSH (green dash–dotted), and ACV_NOSH + IACV_SH (orange dash–dotted).

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Effects of Monsoon Trough Intraseasonal Oscillation on Tropical Cyclogenesis over the Western North Pacific

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  • 1 Center for Monsoon System Research, Institute of Atmospheric Physics, Chinese Academy of Sciences, Beijing, China, and International Pacific Research Center, and School of Ocean and Earth Science and Technology, University of Hawai‘i at Mānoa, Honolulu, Hawaii
  • | 2 International Pacific Research Center, and School of Ocean and Earth Science and Technology, University of Hawai‘i at Mānoa, Honolulu, Hawaii, and CDRC/ESMC, International Laboratory on Climate and Environment Change, Nanjing University of Information Science and Technology, Nanjing, China
  • | 3 Naval Research Laboratory, Monterey, California
  • | 4 Center for Monsoon System Research, Institute of Atmospheric Physics, Chinese Academy of Sciences, Beijing, China
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Abstract

The effects of intraseasonal oscillation (ISO) of the western North Pacific (WNP) monsoon trough on tropical cyclone (TC) formation were investigated using the Advanced Research Weather Research and Forecasting (ARW) Model. A weak vortex was specified initially and inserted into the background fields containing climatological-mean anomalies associated with active and inactive phases of monsoon trough ISOs.

The diagnosis of simulations showed that monsoon trough ISO can modulate TC development through both dynamic and thermodynamic processes. The dynamic impact is attributed to the lower–midtropospheric large-scale vorticity associated with monsoon trough ISO. Interactions between cyclonic vorticity in the lower middle troposphere during the active ISO phase and a vortex lead to the generation of vortex-scale outflow at the midlevel, which promotes the upward penetration of friction-induced ascending motion and thus upward moisture transport. In addition, the low-level convergence associated with active ISO also helps the upward moisture transport. Both processes contribute to stronger diabatic heating and thus promote a positive convection–circulation–moisture feedback. On the other hand, the large-scale flow associated with inactive ISO suppresses upward motion near the core by inducing the midlevel inflow and the divergence forcing within the boundary layer, both inhibiting TC development. The thermodynamic impact comes from greater background specific humidity associated with active ISO that allows a stronger diabatic heating. Experiments that separated the dynamic and thermodynamic impacts of the ISO showed that the thermodynamic anomaly from active ISO contributes more to TC development, while the dynamic anomalies from inactive ISO can inhibit vortex development completely.

School of Ocean and Earth Science and Technology Contribution Number 9130 and International Pacific Research Center Contribution Number 1062.

Corresponding author address: Tim Li, IPRC and SOEST, University of Hawai‘i at Mānoa, 1680 East West Road, POST Bldg. 401, Honolulu, HI 96822. E-mail: timli@hawaii.edu

Abstract

The effects of intraseasonal oscillation (ISO) of the western North Pacific (WNP) monsoon trough on tropical cyclone (TC) formation were investigated using the Advanced Research Weather Research and Forecasting (ARW) Model. A weak vortex was specified initially and inserted into the background fields containing climatological-mean anomalies associated with active and inactive phases of monsoon trough ISOs.

The diagnosis of simulations showed that monsoon trough ISO can modulate TC development through both dynamic and thermodynamic processes. The dynamic impact is attributed to the lower–midtropospheric large-scale vorticity associated with monsoon trough ISO. Interactions between cyclonic vorticity in the lower middle troposphere during the active ISO phase and a vortex lead to the generation of vortex-scale outflow at the midlevel, which promotes the upward penetration of friction-induced ascending motion and thus upward moisture transport. In addition, the low-level convergence associated with active ISO also helps the upward moisture transport. Both processes contribute to stronger diabatic heating and thus promote a positive convection–circulation–moisture feedback. On the other hand, the large-scale flow associated with inactive ISO suppresses upward motion near the core by inducing the midlevel inflow and the divergence forcing within the boundary layer, both inhibiting TC development. The thermodynamic impact comes from greater background specific humidity associated with active ISO that allows a stronger diabatic heating. Experiments that separated the dynamic and thermodynamic impacts of the ISO showed that the thermodynamic anomaly from active ISO contributes more to TC development, while the dynamic anomalies from inactive ISO can inhibit vortex development completely.

School of Ocean and Earth Science and Technology Contribution Number 9130 and International Pacific Research Center Contribution Number 1062.

Corresponding author address: Tim Li, IPRC and SOEST, University of Hawai‘i at Mānoa, 1680 East West Road, POST Bldg. 401, Honolulu, HI 96822. E-mail: timli@hawaii.edu

1. Introduction

Tropical cyclone (TC) genesis is a process through which disorganized convective-scale cloud systems are developed into a self-sustaining cyclonic system with a warm core under favorable large-scale background conditions. Owing to the lack of reliable observational data over open oceans and complex multiscale interactions involved, the understanding of tropical cyclogenesis remains limited.

The western North Pacific (WNP) is the most active region of tropical cyclogenesis, where more than 30% of global TCs form (Neumann 1993). Gray (1968) identified six necessary conditions for tropical cyclogenesis: warm sea surface temperature (SST) with a deep oceanic mixed layer, high midlevel relative humidity, a conditionally unstable atmosphere, low-level positive relative vorticity, weak vertical wind shear, and nonzero Coriolis force. Low-level convergence and preexisting disturbances with convection are also favorable for TC formation (Zehr 1992). Over the WNP, approximately 70%–80% of TCs develop within a monsoon trough, which is characterized by the monsoon westerly to its south, the trade easterly to its north, and abundant precipitation (Chen and Ding 1979; McBride 1995; Molinari and Vollaro 2013).

The intensity of monsoon trough exhibits significant intraseasonal variability with a dominant frequency of 20–80 days [see Li and Wang (2005) for a review]. During boreal summer, the intraseasonal oscillation (ISO) exhibits marked eastward and northward propagations over the tropical Indian Ocean and northward and northwestward propagations over the South China Sea and WNP (Wang and Rui 1990; Wang and Xie 1997; Jiang et al. 2004). The northward and northwestward propagations of the ISO lead to the significant fluctuation of monsoon trough on the intraseasonal time scale over the WNP (Hsu et al. 2011). Liu and Wang (2014) showed that the seasonal-mean state of monsoon trough plays an essential role in sustaining a strong ISO signal over the WNP. Previous studies indicated that the Madden–Julian oscillation [MJO; Madden and Julian (1994)] plays an important role in TC formation and development (e.g., Gray 1979; Nakazawa 1986; Liebmann et al. 1994; Maloney and Hartmann 2000a,b; Hall et al. 2001; Bessafi and Wheeler 2006; Camargo et al. 2009; Ge et al. 2010; Gao and Li 2011; Huang et al. 2011; Li 2012; Li and Zhou 2013). For example, Gray (1979) suggested that during TC seasons, there are typically 1–2-week active periods, followed by 2–3-week periods of quiescence. Liebmann et al. (1994) and Kim et al. (2008) both showed an increase (decrease) of TC formation during the active (inactive) MJO phase over the WNP. Li and Zhou (2013) indicated that the significant change of TC frequency on the intraseasonal time scale is associated with the strengthening and weakening of monsoon trough over the WNP. It has been shown that the dynamic impact of circulation associated with MJO on TCs is through barotropic eddy kinetic energy conversion by which disturbances obtain energy from the large-scale environment flow (Maloney and Hartmann 2001; Hsu et al. 2011; Hsu and Li 2011; Cao et al. 2012, 2013). The amplitude and spatial pattern of the barotropic energy conversion, however, depend on the rotational and divergent components of the flow within the MJO. While the rotational flow contribution is about 50% larger, it primarily affects TCs in the northern part of the WNP basin, and the divergent flow contribution is mainly confined to the southern part of the basin (Hsu et al. 2011). In addition to barotropic energy conversion, MJO may modulate tropical cyclogenesis through the change of background low-level convergence, low-level vorticity, vertical shear, and humidity (Maloney and Hartmann 2000a, 2001; Camargo et al. 2009).

Mao and Wu (2010) and Cao et al. (2012) examined the circulation patterns associated with the active and inactive phases of monsoon trough ISO over the WNP. They noted that an active monsoon trough ISO phase is associated with a cyclonic (anticyclonic) circulation anomaly in the lower (upper) troposphere, enhanced midtropospheric ascending motion, and higher relative humidity in the middle troposphere. Many previous studies have suggested that such large-scale background conditions during active ISO phase favor TC development. However, the relative contributions of the dynamic and thermodynamic processes to TC formation during this period have not yet been established. This work attempts to address the proposed question by employing idealized numerical model simulation.

The rest of this paper is organized as follows. The model configuration, large-scale anomalies of monsoon trough ISO, initial vortex, and experimental design are illustrated in section 2. In section 3, the general vortex structure and their time evolutions are examined. The modulations and mechanisms of monsoon trough ISO on TC formation are examined in section 4. In section 5, the relative roles of ISO dynamic and thermodynamic impacts in TC formation are further examined with additional experiments. Finally, major findings of the study are summarized and some discussions are given in section 6.

2. The model and experiment design

a. Model configuration

The model used in this study is the nonhydrostatic Advanced Research Weather Research and Forecasting (ARW) Model (version 3.3) developed by the National Center for Atmospheric Research (Skamarock et al. 2008). The model has two inner nests with two-way interaction. The mesh sizes in the three domains are 241 × 241, 241 × 241, and 481 × 481 with horizontal grid sizes of 27, 9, and 3 km, respectively. The Kain–Fritch convective scheme is applied to the two outermost meshes (Kain and Fritsch 1993) and an explicit microphysics scheme (Lin et al. 1983) is used in all meshes. Other model physics include the Yonsei University (YSU) planetary boundary layer (PBL) scheme, thermal diffusion land surface scheme, and Monin–Obukhov surface-layer scheme (Hong et al. 2006). A fixed lateral boundary condition is used for the outermost domain, in which the tendency of prognostic variables in the lateral boundary is equal to zero. The model is set on a beta plane with a weak vortex placed at 20°N initially. In the control experiment, the background environment flow is quiescent and the SST is a constant of 29°C. This specified SST value is consistent with the region of the WNP under consideration during June–September. The relative humidity and other thermodynamic variables are horizontally homogeneous and have the vertical profile of typical January-mean observations at Willis Island, northeast of Australia (Holland 1997), which is similar to the July-mean observations over the WNP. The details for the model configuration are given in Table 1.

Table 1.

The model configuration.

Table 1.

b. Large-scale anomalies of monsoon trough ISO

To examine how monsoon trough ISO may modulate vortex development, the large-scale anomaly fields associated with the active and inactive phases of monsoon trough ISO derived from the National Centers for Environmental Prediction (NCEP) reanalysis data (Kalnay et al. 1996) were added to the mean background for the ISO experiments. It is worth mentioning that the ISO patterns derived from the NCEP reanalysis are generally similar to those derived from other reanalysis products such as the Japanese 25-year Reanalysis Project (JRA-25). First, the daily satellite outgoing longwave radiation (OLR) dataset from the National Oceanic and Atmosphere Administration (NOAA) (Liebmann and Smith 1996) was used to compute an intensity index of monsoon trough ISO over the WNP. Then the NCEP reanalysis data covering a period from 1979 to 1998 during the June–September WNP TC seasons were used to compute monsoon trough ISO anomaly. The annual cycle was removed from the long-term daily mean for the period. A Lanczos bandpass filter was performed to extract the 20–80-day oscillation signals (Duchon 1979). Note that both datasets had a horizontal resolution of 2.5° in latitude and longitude.

The distribution of 20–80-day-filtered OLR anomaly variance from June to September shows that a maximum ISO intensity center appears along the monsoon trough (Figs. 1a,b). According to the ISO maximum variance of OLR anomaly, a box-averaged (12°–22°N, 123°–133°E) OLR index was used as a reference time series to select active and inactive ISO events. The active and inactive phases of monsoon trough ISO were defined according to the normalized box-averaged OLR index (with its absolute amplitude greater than one standard deviation). The positive (negative) values represented the suppressed (enhanced) convection in the monsoon trough during the inactive (active) phase of ISO. The composite technique was then applied to obtain anomalous circulation fields during the active and inactive phases of ISO. If we expanded the OLR index to a larger domain (12°–22°N, 123°–143°E), the vertical distributions of dynamic and thermodynamic variables during the active and inactive phases were almost identical (figure not shown). A similar result could be derived when we expanded the box to 12°–22°N, 113°–143°E. This indicated that the active and inactive phases of monsoon trough ISO were not sensitive to the choice of relatively small or large domain. The active ISO phase is characterized by a large-scale cyclonic circulation anomaly in the low levels. The width of the cyclonic system is approximately 4000 km (Fig. 1c). A nearly mirror image pattern with an anticyclonic circulation appears in the inactive ISO phase (Fig. 1d).

Fig. 1.
Fig. 1.

(a) Climatological-mean winds at 850 hPa and (b) variance of 20–80-day-filtered OLR anomalies during June–September for the period 1979–98. The dashed box in (a) and (b) represents the major monsoon trough region (12°–22°N, 123°–133°E) for an intensity index, and areas more than 400 (W m−2)2 in (b) are shaded. The composite wind at 850 hPa during (c) the active phase and (d) the inactive phase. The dashed box in (c) and (d) represents the area (10°–30°N, 112°–132°E) in Fig. 2.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

Figure 2 shows the vertical profiles of area-averaged (10°–30°N, 112°–132°E) tangential wind, temperature, divergence, and specific humidity centered in the cyclonic and anticyclonic circulation anomalies for the active and inactive phases of the ISO. In the active phase, the positive cyclonic tangential wind anomalies extend from the surface to 250 hPa, with a maximum magnitude near 700 hPa, changing into anticyclonic wind anomalies above 250 hPa (Fig. 2a, solid line). Positive temperature anomalies appear throughout the troposphere, with a maximum center at 200 hPa (Fig. 2b). Convergence dominates below 700 hPa and divergence prevails above 300 hPa (Fig. 2c), consistent with the tangential wind profile. Positive specific humidity anomalies appear throughout the whole troposphere, with a maximum magnitude (0.5 g kg−1) in PBL (Fig. 2d). The vertical profiles of monsoon trough ISO during the inactive phase are opposite to those for the active phase with slightly different magnitudes (dashed lines in Fig. 2).

Fig. 2.
Fig. 2.

The vertical profiles of area-averaged 20° × 20° (a) tangential wind, (b) temperature, (c) divergence, and (d) specific humidity during the active (solid) and inactive (dashed) phases of monsoon trough ISO centered in the circulation anomaly. The vertical dashed line indicates the zero value of each parameter.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

c. Initial vortex

An axisymmetric weak vortex is placed at the center of the domain for the control run and at the large-scale ISO circulation centers for the ISO experiments. The tangential wind of the initial balanced vortex has the following radial and vertical profiles:
e1
where r is the radial distance from the vortex center, is the radius of maximum tangential wind, is the maximum tangential wind at the radius of , and is vertical sigma level. The maximum tangential wind of the initial vortex is 8 m s−1 at a radius of 150 km at the surface. The wind speed gradually decreases upward. This initial bogus vortex has a minimum sea level pressure (MSLP) of 1010 hPa. Given Eq. (1), the mass and thermodynamic fields are derived based on a nonlinear balance equation so that the initial vortex satisfies both the hydrostatic and gradient wind balances (Wang 1995, 2001). In this study, we specify a bottom vortex rather than a midlevel vortex, and as a result the initial vortex has a warm-core structure. It has been shown (e.g., Ge et al. 2013) that a bottom vortex is more efficient in tropical cyclogenesis than a midlevel vortex.

d. Experimental design

Three sets of experiments were designed first, given in Table 2. In the control experiment, an initial bogus vortex was placed in a resting environment (hereafter CTL). In the second group of experiments, the active ISO anomalies that include all the dynamic and thermodynamic fields were added to the background state. In this active ISO group, two experiments were conducted; one with an initial weak vortex and the other one without a vortex, identified as ACV and AC, respectively. The difference between the results of these two experiments will represent the impact of ISO on vortex evolution. In the third group of experiments, the inactive ISO anomalies were added to the background fields. The experiments with and without a vortex were identified as IACV and IAC, respectively.

Table 2.

Numerical experiments.

Table 2.

To understand more clearly the dynamic and thermodynamic effects associated with the ISO, two additional groups of experiments were designed (see Table 3). In the first group of the experiments, the dynamic variables including anomalies of zonal wind u, meridional wind υ, surface pressure (ps), geopotential height (hgt), and temperature T from the active ISO phase composite were added to the background fields, while the specific humidity anomaly associated with the active ISO was set to zero. In the second group of experiments, all dynamic anomalies associated with the active phase of the ISO were set to zero, while the specific humidity anomaly associated with the active phase of the ISO was added to the background field. Both groups of experiments contained runs with and without the initial vortex. These additional sets of experiments were listed in Table 3.

Table 3.

Sensitivity experiments.

Table 3.

Note that the experiment name prefix “AC” or “IAC” denotes active or inactive phase of monsoon trough ISO, “V” denotes the presence of an initial bogus vortex, “SH” denotes the experiment with prescribed ISO moisture field but not the dynamic fields, and “NOSH” denotes the experiment with prescribed ISO dynamic fields but not the moisture field.

3. Vortex evolutions during active and inactive ISO phases

Because ACV and IACV included both the vortex perturbation and large-scale ISO fields, to examine “pure” vortex evolutions, the ISO simulated fields from AC and IAC were subtracted from the ACV and IACV fields, respectively. The resulting fields consist of two parts. The first part represents the pure vortex evolution under the resting environment (similar to one in CTL), and the second part represents additional vortex evolution induced by the interaction of the vortex with the large-scale ISO field.

Figure 3 shows the simulated wind fields at 850 hPa at t = 0, 24, and 48 h in CTL, ACV, and IACV with the large-scale ISO field removed. The initial vortices are identical in these three cases (Figs. 3a,d,g). With the beta effect, a wavenumber-1 beta gyre is induced, causing a northwest–southeast orientation of the vortices as shown at t = 24 and 48 h (Figs. 3b,c,e,f,h,i). Among the three experiments, the vortex in ACV is most symmetric, while the one in IACV is least symmetric. This reflects favorable (unfavorable) conditions in ACV (IACV).

Fig. 3.
Fig. 3.

Wind fields (vectors) and total wind speed (shaded) at 850 hPa in (a)–(c) CTL, (d)–(f) ACV relative to AC, and (g)–(i) IACV relative to IAC at t = (left) 0, (center) 24, and (right) 48 h from 27-km simulation data.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

Because of the limitation of the domain, our integration is carried out to 120 h with the display up to 108 h only so that the TCs are not affected by the lateral boundaries yet. Figure 4 shows the time evolutions of the MSLP and the maximum azimuthal-mean wind (MAMW) at the 10-m height in CTL, ACV, and IACV. In the first 48 h the vortices rarely develop for all three cases, followed by three different evolution sequences. A strong TC develops at t = 108 h in ACV, whereas the vortex fails to grow throughout the whole 108-h simulation in IACV. In the meantime, the vortex in CTL without the ISO background flow grows more slowly compared to ACV. When the MAMW at the 10 m is greater than 15 m s−1, we define this time as the tropical cyclogenesis time. Based on this definition, tropical cyclogenesis appears at t = 99 h in CTL and at t = 72 h in ACV, respectively.

Fig. 4.
Fig. 4.

Time evolution of (a) the MSLP and (b) the MAMW speed at 10 m for CTL (black solid), ACV (red dashed), and IACV (blue dotted). The abscissa represents time (h) and the ordinate corresponds to the intensity.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

These experiments indicate that the large-scale dynamic and thermodynamic conditions associated with the active (inactive) phase of monsoon trough ISO provide a more (less) favorable environment for TC formation. This result is consistent with many previous observational analyses (e.g., Liebmann et al. 1994; Maloney and Hartmann 2001). Next, we investigate specific processes through which the ISO dynamic and thermodynamic variables influence vortex development.

4. Mechanisms for ISO impact on vortex development

Figure 5 shows the radial profiles of the vortex azimuthal-mean radial wind, tangential wind, and geopotential at 1000 hPa in the first 7 h of the ACV and IACV experiments at a 2-h interval. While the tangential wind components are similar in this first 7-h display (Figs. 5e–h), the radial wind and the geopotential exhibit greater differences between the active and inactive ISO experiments, especially by t = 7 h (Figs. 5d,l). The radial wind is predominantly inflow in the IACV case, but an outflow develops in the inner core in the ACV case (Fig. 5d). This suggests that the ISO impact is mainly through the radial wind component (as radial pressure gradient force is associated with the radial wind acceleration). We will diagnose the radial wind momentum budget. The radial wind momentum equation in cylindrical and pressure coordinates can be written as follows:
e2
Assume that total wind is the sum of background flow and vortex perturbation (i.e., ), where an overbar denotes the ISO flow and a prime denotes the vortex perturbation. Assuming that the ISO background flows satisfy Eq. (2), the difference between ACV (IACV) and AC (IAC) would represent the perturbation (vortex) evolution. Thus, the perturbation momentum equation may be written as
e3
where is azimuthal angle; r is radius; p is pressure; , , are radial, tangential, and vertical wind speeds of the background flow, respectively; is the Coriolis parameter; , , , and are radial, tangential, vertical winds, and the geopotential of the perturbation (vortex), respectively; and is horizontal and vertical diffusion in the radial direction. In Fig. 5, an azimuthal mean for Eq. (3) was taken to get the radial distribution of the vortex.
Fig. 5.
Fig. 5.

The radial distribution of the (a)–(d) azimuthal-mean perturbation radial wind, (e)–(h) tangential wind, and (i)–(l) geopotential at 1000 hPa from t = 1 to 7 h with an interval of 2 h for ACV relative to AC (black) and IACV relative to IAC (red).

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

While friction-induced near-surface inflow occurs under both the active and inactive ISO flows, a major difference appears in the middle troposphere compared to CTL (Fig. 6). At t = 1 h, there is clearly an outflow in the active phase (Fig. 6a) but an inflow in the inactive phase in the layer between 600 and 800 hPa (Fig. 6e). Given an initially symmetric and gradient wind balanced vortex, and are equal to zero and is satisfied. Therefore, two terms in perturbation momentum equation [Eq. (3)] that can drive radial wind acceleration in the free atmosphere (where friction can be neglected) are and . Equation (3) at initial time is then reduced to . Our calculation shows that the second term is very small and the major difference between the active and inactive phases lies in the first term (). This term is positive with a maximum around 700 hPa during an active phase because both the ISO and vortex flows are cyclonic and the maximum ISO circulation is around 700 hPa (Figs. 7a and 2a). This term is negative in an inactive phase as the ISO flow is anticyclonic (Figs. 7b and 2a). As a consequence of the interaction between the ISO and vortex flows, a radial outflow (inflow) is generated initially below 600 hPa and maximized around 700 hPa during the active (inactive) phase (Figs. 6a,e).

Fig. 6.
Fig. 6.

The vertical–radial cross section of azimuthal-mean perturbation radial wind (m s−1) from t = 1 to 7 h with an interval of 2 h for (a)–(d) ACV relative to AC and for (e)–(h) IACV relative to IAC and (i)–(l) CTL.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

Fig. 7.
Fig. 7.

The vertical–radial cross section of azimuthal-mean (m s−2) at t = 1 h between the (a) active phase and (b) inactive phase.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

What is the impact of the distinctive midlevel outflow and inflow on subsequent vortex development? Recall that the near-surface inflow begins to develop in both active and inactive phases because the surface friction decreases the tangential wind and leads to a gradient wind imbalance. For the mass continuity, outflows are induced immediately above the inflow layer. During the active phase, the near-surface inflow is well connected with the midlevel outflow (Fig. 6a). This differs from the inactive phase case, in which only a very narrow outflow layer occurs between the inflows near the surface and at 700 hPa (Fig. 6e). The midlevel inflow also causes the outflow above 500 hPa owing to the mass continuity in the inactive phase (Fig. 6e). Therefore, the different background ISO flows lead to distinctive radial wind vertical profiles above the surface inflow layer in the early stage. As the boundary layer inflow enhances, the outflow layer above it intensifies and thickens and pushes the inward-flow layer upward for both the active and inactive cases by t = 3 h (Figs. 6b,f). The inward-flow layer keeps steady with a maximum around 500 hPa after t = 3 h in the inactive phase, which, through induced subsidence in the vortex-core region below 500 hPa, has a lasting effect on the vortex development beyond t = 3 h. While radial wind profiles look similar in general below 700 hPa for the active and inactive ISO cases at t = 5 h, they show marked differences at t = 7 h, especially within the boundary layer where there is a convergence in the ACV but not in the IACV (Figs. 6d,h). Meantime, a comparison of Figs. 6d and 6h shows a clear difference in the depth of the secondary circulation at t = 7 h; in the active phase, “in–up–out” secondary circulation extends from the surface to around 700 hPa, whereas in the inactive phase it is confined below 900 hPa. This difference warrants a more detailed analysis.

In Fig. 8, the simulated fields are shown for the ACV case with a smaller time interval. Between t = 6 and 6.5 h (Figs. 8b,c), the radial wind develops an outward flow near the core within the boundary layer (below 900 hPa), which is accompanied by strengthened, upward-penetrated ascending motion and diabatic heating. It is the combination of three factors—strengthened frictional inflow, the midlevel outflow induced by the ISO active field, and the low-level convergence of the ISO active field (solid line in Fig. 2c)—that contribute the upward development of perturbation vertical motion and associated heating. The heating lowers local surface pressure, promoting a stronger convergence near the surface at 135 km, as revealed by the radial wind and the convergence–divergence fields associated with the vortex. This signal is weak at t = 6 h and becomes much stronger at t = 6.5 h (see Figs. 8b,c and 8e,f). The vertical velocity and diabatic heating are confined within the boundary layer at t = 5.5 h (Figs. 8g,j). With a weak convergence forming (Fig. 8e), the vertical velocity starts protruding upward at t = 6 h (Fig. 8h). With a strong low-level convergence developing (Fig. 8f), deep convection finally is triggered by t = 6.5 h (Fig. 8i). The greater ascending motion and upward moisture transport (figure not shown) contribute to a greater convective heating (Fig. 8l). The heating can further strengthen the low-level inflow by changing the temperature and surface pressure. Through this positive feedback process, the vortex develops.

Fig. 8.
Fig. 8.

The vertical–radial cross section of (a)–(c) azimuthal-mean perturbation radial wind (m s−1), (d)–(f) divergence (10−5 s−1), (g)–(i) vertical velocity (10−3 m s−1), and (j)–(l) diabatic heating (10−4 K s−1) at (left) t = 5.5, (center) 6, and (right) 6.5 h for ACV relative to AC.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

For the inactive ISO case, the ISO flow imposes an inward flow in the middle troposphere in the early stage (Figs. 6e–g), inducing anomalous subsidence below and suppressing the vortex secondary circulation. Meanwhile, the inactive ISO anomaly imposes a divergence on the vortex in the low level (Fig. 2c, dashed line). Therefore, vortex ascending motion associated with frictional convergence is confined near the top of the boundary layer and cannot develop into a higher level, which hinders all the subsequent critical elements such as boundary layer outflow inside of the maximum ascending motion and vertically penetrated diabatic heating as seen in the active ISO phase (Figs. 9d–f). The vortex vertical motion and diabatic heating remain very shallow (Figs. 9g–l). Thus, a TC could not form in the environment of an inactive ISO.

Fig. 9.
Fig. 9.

As in Fig. 8, but for IACV relative to IAC.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

Figure 10 shows the time evolutions of diabatic heating and vertical motion profiles of the vortices in the active and inactive ISO cases during the first 10 h. In the active ISO case, initial heating is very shallow between t = 4 and 6 h, and it then penetrates into the upper troposphere beyond t = 6 h (Fig. 10a). The evolution of the diabatic heating corresponds well with the vertical motion (Fig. 10c). The triggering of strong diabatic heating after t = 6 h depicts the development of the boundary layer convergence and vertical motion discussed and shown in Figs. 8f and 8i at t = 6.5 h. In the inactive ISO case, the heating and ascending motion remain very shallow (confined below 850 hPa) during the same time (Figs. 10b,d). This lack of deep convection is the reason that the vortex never develops into a TC.

Fig. 10.
Fig. 10.

Vertical–time cross section of (a),(b) perturbation diabatic heating (10−4 K s−1) and (c),(d) perturbation vertical velocity (10−2 m s−1) averaged over 420 km × 420 km centered on the TC center in the first 10 h for (left) ACV relative to AC and (right) IACV relative to IAC.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

The discussion above is focused on the dynamic impact of the ISO. In addition to this dynamic impact, the active phase of the ISO may also strengthen the vortex development through the thermodynamic (moisture) effect. Recall that the background moisture is larger during the active ISO phase while it is smaller during the inactive ISO phase (Fig. 2d). Such a background moisture difference can affect the vortex development. This is because even given the same background flows, the enhanced background moisture during the active phase would induce more convective activities and promote a greater diabatic heating, thus leading to a greater intensification rate of vortex.

The discussion above mainly focuses on what causes the difference of vortices in the early development stage during the active and inactive phases. To demonstrate that the difference in the initial 10 h is critical in affecting the subsequent vortex development, Fig. 11 shows time evolutions of area-averaged radial wind, tangential wind, diabatic heating, and vertical motion from t = 12 to 60 h in ACV and IACV runs. As we know, TC-like vortex development experiences an oscillation development (Li et al. 2006). Thus, to compare the overall strength of two vortices, it is necessary to use area- and time-averaged variables rather than those variables at a single time and a single spatial point. For this reason, we plotted 180-km radial-averaged, 24-h-running-mean symmetric components of these fields. Figure 11 clearly illustrates, from all four variables selected, that the difference of vortex intensity in the early stage does impact the subsequent evolution and final TC intensity. Physically, it is reasonable that more vertically penetrated ascending motion in ACV during the initial 10 h would favor greater diabatic heating, which would lead to further development of PBL inflow and ascending motion.

Fig. 11.
Fig. 11.

Time (24-h running mean) evolutions of radial-mean (0–180 km) (a) radial wind at 10 m, (b) tangential wind at 10 m, (c) vertically integrated (1000–200 hPa) diabatic heating, and (d) vertically integrated (1000–200 hPa) vertical velocity for ACV relative to AC (red) and IACV relative to IAC (blue) cases.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

In addition to the above area-averaged quantities, we also plotted the vertical–radial cross section of azimuthal-mean vertical velocity and diabatic heating for the ACV and IACV, averaged at 0–24, 24–48, and 48–72 h (Fig. 12). It is clear that the difference of the vertical velocity and diabatic heating between ACV and IACV in the early stage can indeed make an impact to the subsequent vortex development.

Fig. 12.
Fig. 12.

The vertical–radial cross section of azimuthal-mean (a)–(f) vertical velocity (10−2 m s−1) and (g)–(l) diabatic heating (10−4 K s−1) fields for (a)–(c),(g)–(i) ACV relative to AC and (d)–(f),(j)–(l) IACV relative to IAC cases, averaged at 0–24, 24–48, and 48–72 h.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

In the next section, we further separate the dynamic and thermodynamic factors from the ISO experiments to reveal their relative importance.

5. Relative roles of the ISO dynamic and thermodynamic impacts

Table 3 describes two sets of experiments designed to investigate separately the dynamic and thermodynamic (mainly moisture) impacts of the active ISO on vortex development. In the ACV_SH experiment, only the specific humidity anomaly from the active ISO is added and the ACV_NOSH case contains the dynamic anomalies from the active ISO only. Figure 13 shows the time evolutions of simulated MSLP and MAMW in CTL, ACV, ACV_SH, and ACV_NOSH for the ISO active experiments. Compared to CTL, both the dynamic and thermodynamic parameters associated with the active ISO phase contribute positively to vortex development. In addition, the thermodynamic component of monsoon trough ISO has a greater contribution to vortex development than the dynamic components.

Fig. 13.
Fig. 13.

As in Fig. 4, but for CTL (black solid), ACV (red dashed), ACV_SH (blue dotted), and ACV_NOSH (green dash–dotted).

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

The relative magnitude of the dynamic and thermodynamic contributions can be estimated by the MSLP (MAMW) differences between ACV_SH and CTL and between ACV_NOSH and CTL. Table 4 lists the MSLP and MAMW differences between ACV_SH and CTL and between ACV_NOSH and CTL for some instantaneous values from t = 84 to 108 h and the average from t = 84 to 108 h. The averaged benefit from the enhanced moisture of the active ISO is approximately double the benefit from the dynamic fields, as seen from both the MSLP and MAMW fields. The result indicates that the thermodynamic factor associated with active monsoon trough ISO plays a greater role in affecting the TC development (Nolan 2007; Wang 2014), while the dynamic factors are also important. This is based on the climatological-mean variables retrieved from the reanalysis fields. For an individual case, the relative importance of the thermodynamic versus the dynamic controls may be different depending on the relative magnitude of these variables.

Table 4.

The TC instantaneous intensity differences and averaged differences from t = 84 to 108 h in terms of the MSLP (hPa) and the MAMW (m s−1) between ACV_SH and CTL and between ACV_NOSH and CTL.

Table 4.

To examine in more detail the relative impact of the dynamic and thermodynamic components, the vertical–radial cross sections of relevant fields at t = 5, 7, and 9 h for CTL, ACV_SH, and ACV_NOSH are plotted in Figs. 1416, respectively. In the CTL run, a weak convergence develops at t = 7 h and it amplifies by t = 9 h, even without the help of large-scale convergence from the ISO (Figs. 14e,f). Not surprisingly, the radial wind vertical profile in ACV_NOSH (Figs. 15a,b) is, in general, similar to that in ACV (Figs. 6c,d), while the radial wind vertical profile in ACV_SH (Figs. 16a,b) is similar to that in CTL (Figs. 14a,b) at t = 5 and 7 h. The corresponding divergence fields in ACV_SH (Figs. 16d,e) are similar to the CTL (Figs. 14d,e) at t = 5 and 7 h. However, difference arises for the vertical velocity and associated diabatic heating at t = 7 h (Figs. 14h,k and 16h,k). With larger moisture field in the ACV_SH, deep convection is triggered earlier and much deeper diabatic heating is generated by t = 9 h (Fig. 16l). In the meantime, convective activity remains shallow in the CTL (Fig. 14l) and the TC development is slower (Fig. 13).

Fig. 14.
Fig. 14.

The vertical–radial cross section of azimuthal-mean (a)–(c) perturbation radial wind (m s−1), (d)–(f) divergence (10−5 s−1), (g)–(i) vertical velocity (10−3 m s−1), and (j)–(l) diabatic heating (10−4 K s−1) at (left) t = 5, (center) 7, and (right) 9 h for CTL run.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

Fig. 15.
Fig. 15.

As in Fig. 14, but for ACV_NOSH relative to AC_NOSH.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

Fig. 16.
Fig. 16.

As in Fig. 14, but for ACV_SH relative to AC_SH.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

The dynamic control of the active ISO without the accompanying moisture anomaly (ACV_NOSH) is examined by comparing Figs. 14 and 15. The diabatic heating remains very shallow in the CTL run by t = 9 h (Fig. 14l). On the other hand, with the help of convergence from the ISO dynamics, the convergence for the vortex is stronger and allows a stronger vertical velocity that penetrates upward (Figs. 15h,i). The accompanying diabatic heating is deeper than that in the CTL run (Figs. 15l and 14l).

Comparing the dynamic (ACV_NOSH) and the moisture (ACV_SH) impacts from the active ISO, the low-level convergence is greater in the dynamic impact case so that the vertical motion and the accompanying diabatic heating is much stronger in the early stage shown at t = 5 and 7 h (Figs. 15g,h,j,k and 16g,h,j,k). However, because of less moisture in ACV_NOSH, the system cannot support as strong and deep convection as in the ACV_SH case (as illustrated at t = 9 h). Therefore, the development of TC vortex is weaker in ACV_NOSH than in ACV_SH.

The distinction among the three experiments can be clearly illustrated in the time evolutions of the area-averaged vertical profiles of the vertical velocity and diabatic heating (Fig. 17). In comparison with CTL, a slightly earlier development of ascending motion in the low levels is observed in ACV_NOSH at t = 3 h (Fig. 17c). The overall pattern of the development of the vertical motion is similar between CTL and ACV_NOSH except a greater magnitude is observed in ACV_NOSH (Figs. 17a,c). This is consistent with a greater radial inflow near the surface in ACV_NOSH. In the absence of the ISO moisture effect, the dynamic factors alone help the earlier development of the vertical motion through the low-level ISO convergence forcing as discussed in Fig. 15. In contrast, owing to the lack of the dynamic forcing from the ISO, the difference of vortex vertical velocity between ACV_SH and CTL is very small within the first 6 h (Figs. 17a,e). However, with more abundant moisture associated with the active ISO in the large-scale environment, more vigorous convection allows to form. The magnitudes of both the vertical velocity and diabatic heating fields start to deviate from each other between ACV_SH and CTL (Figs. 17a,b,e,f). This is because the enhanced background moisture can support stronger convection and vertical motion with strengthened diabatic heating, which can promote a greater positive feedback loop through the change of the perturbation dynamic fields. Note that the evolutions of diabatic heating in these three experiments all show a transition from a shallow heating mode to a deep heating mode, consistent with the two-stage conceptual model of transition from cumulus congestus to deep convection proposed by Wang (2014).

Fig. 17.
Fig. 17.

The vertical–radial cross section of (a),(c),(e) perturbation vertical velocity and (b),(d),(f) perturbation diabatic heating averaged over 420 km × 420 km centered on the TC center in the first 10 h for (a),(b) CTL, (c),(d) ACV_NOSH relative to AC_NOSH, and (e),(f) ACV_SH relative to AC_SH.

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

The comparison of Figs. 17c and 17d with Figs. 17e and 17f shows that the dynamic factors of the active ISO exert a greater impact on the vortex development during the initial 6 h, whereas the thermodynamic effect of the active ISO becomes more important in the later development stage.

We also examine separately the dynamic and thermodynamic effects from the inactive ISO phase that hinder the vortex development. As labeled for the active ISO experiments, the IACV_NOSH includes the dynamic anomalies from the inactive ISO while the IACV_SH includes the thermodynamic anomaly. Figure 18 depicts the time evolution of the four experiments: CTL, IACV, IACV_SH, and IACV_NOSH. Recall that the impact of inactive ISO to TC development is detrimental; the evolution of the vortex in IACV_NOSH is almost the same as in IACV, which contains both the dynamic and thermodynamic anomalies from the inactive ISO. Without the dynamic impact, there is a very small development of the vortex in IAC_SH. This set of experiments for the inactive ISO phase shows that the dynamic impact is greater than the thermodynamic impact in inhibiting the vortex development, albeit insignificant.

Fig. 18.
Fig. 18.

As in Fig. 4, but for CTL (black solid), IACV (red dashed), IACV_SH (blue dotted), and IACV_NOSH (green dash–dotted).

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

Our last set of experiments combines the thermodynamic and dynamic anomalies of active and inactive ISO in a mixed manner. The “ACV_SH + IACV_NOSH” experiment has the dynamic anomalies coming from the inactive ISO and the thermodynamic (moisture) anomaly from the active ISO. The “ACV_NOSH + IACV_SH” experiment has the dynamic anomalies coming from the active ISO and the thermodynamic (moisture) anomaly from the inactive ISO. Figure 19 depicts the time evolution of these two experiments plus CTL, ACV, and IACV as shown in Fig. 4. When the dynamics are coming from the inactive ISO phase, even with the enhanced moisture field from the active ISO phase, the vortex fails to develop. When the dynamic anomalies from the active ISO are included, reduced moisture from the inactive ISO (Fig. 2d) still allows some vortex development. The vortex intensity in “ACV_NOSH + IACV_SH” is close to the one in ACV_NOSH (Fig. 13), suggesting that the reduction of moisture from the inactive ISO only has a minor effect. On the other hand, both “ACV_SH + IACV_NOSH” and IACV_NOSH have no vortex developments, indicating the inhibiting mechanism associated with IACV_NOSH could not be overcome by increasing moisture from the active ISO anomaly. Based on all these experiments, we conclude that the dynamic fields play an upper hand in controlling the TC development.

Fig. 19.
Fig. 19.

As in Fig. 4, but for CTL (black solid), ACV (red dashed), IACV (blue dotted), ACV_SH + IACV_NOSH (green dash–dotted), and ACV_NOSH + IACV_SH (orange dash–dotted).

Citation: Journal of the Atmospheric Sciences 71, 12; 10.1175/JAS-D-13-0407.1

The finding of the relative importance of the dynamic and thermodynamic impacts from the active and inactive ISO phases from our designed experiments may be different in individual case as the background anomalies taken from the active and inactive ISO phases are from the climatological-mean fields.

6. Summary and discussion

Many previous studies have indicated that the environment associated with the MJO active phase provides favorable large-scale background conditions for TC formation, but specific processes through which the MJO characteristics influence the tropical cyclogenesis are unclear. In this study, we design a number of idealized numerical experiments that include the climatological-mean states of the active and inactive phases of monsoon trough ISO from a long period of global reanalysis fields to investigate how they impact the vortex development. We also separately identify the dynamic and thermodynamic impacts of monsoon trough ISO during active and inactive phases.

Our simulations show that the dynamic impact of large-scale ISO circulation on TC development is through the following processes. Interactions between ISO large-scale cyclonic (anticyclonic) flow and vortex tangential wind cause the formation of radial outflow (inflow) in the midlevel during the active (inactive) phase of the ISO. In the presence of active ISO anomalies, the midlevel outflow, along with the low-level ISO convergence, promotes the upward development of perturbation ascending motion induced by friction-induced boundary layer inflow. This strengthens upward moisture transport and induces strong convective heating. The enhanced heating can further affect local temperature and surface pressure, inducing even stronger “in–up–out” secondary circulation. Through this positive feedback, the vortex develops quickly. The numerical experiments demonstrate that in the early stage, the aforementioned dynamic processes play a critical role in supporting vortex development during the active ISO phase. Past the early stage, more abundant moisture coming from the active ISO phase further contributes to the vigorous convection. While in the presence of inactive ISO anomalies, an inflow layer is induced in the middle troposphere, which, along with the large-scale divergent forcing from the inactive ISO, inhibits the upward development of friction-induced ascending motion near the core region. As a result, the vortex cannot develop into a TC.

Further experiments were conducted to examine the relative contribution of dynamic and thermodynamic factors with different and mixed characteristics from the active and inactive ISO phases. It can be seen that the thermodynamic (moisture) factor contributes more to the TC development than the dynamic factors during the active ISO phase. When the dynamic fields are from the ISO inactive phase, the environment becomes hostile so that even with the enhanced moisture field from the active ISO phase, the vortex fails to develop. The overall control of the dynamic fields is in the early stage of the development and the moisture component is responsible for stronger and deeper convection development in the later stage. Based on the long-term ISO composite fields, our conclusion is that the dynamic control is more important in TC development than the thermodynamic control.

In the current study, we focused on the vortex intensification rate difference under the active and inactive ISO phases. In future work, we will further examine the vortex size difference. Observations show that in addition to pronounced ISO, the monsoon trough also exhibits a strong interannual fluctuation. To what extent does the interannual variation of monsoon trough modulate TC development and what are differences of the large-scale controls between monsoon trough interannual variation and ISO? These issues will be addressed elsewhere.

Acknowledgments

X.C. would like to thank Prof. Xuyang Ge for his assistance in numerical model simulation. The authors thank the three anonymous reviewers for helpful comments. This study was supported by China National 973 Project 2015CB453200; National Natural Science Foundation of China (Grants 41475084 and 41275001); Office of Naval Research Grant N00014-1210450; National Science Foundation Grant AGS-1106536; and the International Pacific Research Center (IPRC), which is sponsored by the Japan Agency for Marine-Earth Science and Technology (JAMSTEC).

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